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JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 7
PAGES 1301–1320
2001
High fO2 During Sillimanite Zone
Metamorphism of Part of the Barrovian
Type Locality, Glen Clova, Scotland
JAY J. AGUE∗, ETHAN F. BAXTER† AND JAMES O. ECKERT JR
DEPARTMENT OF GEOLOGY AND GEOPHYSICS, YALE UNIVERSITY, PO BOX 208109, NEW HAVEN, CT 06520-8109, USA
RECEIVED NOVEMBER 8, 1999; REVISED TYPESCRIPT ACCEPTED DECEMBER 5, 2000
The redox state of sillimanite zone (650–700°C, 5–6 kbar)
metasediments of the Barrovian type area, Scotland, was investigated
using estimates of metamorphic oxygen fugacity (fO2 ), sulfur fugacity
(fS2 ), and fluid chemistry based on new determinations of mineral
and rock compositions from 33 samples. A total of 94% of the
samples lack graphite, contain both ilmenite–hematite solid solutions
(RHOMOX) and magnetite, and had metamorphic fO2 about 2
log10 units above the quartz–fayalite–magnetite (QFM) buffer. The
regional variation in metamorphic fO2 for these rocks was minimal,
about ±0·3 log10 units, reflecting either a protolith that was
homogeneous with respect to redox state, or an initially variable
protolith whose redox state was homogenized by metamorphic
fluid–rock interaction. RHOMOX inclusions in garnet porphyroblasts that become richer in ilmenite from the interior to the
edge of the host porphyroblast suggest that at least some synmetamorphic reduction of rock occurred. Significant variations in
bulk-rock oxidation ratio (OR) that are probably inherited from
sedimentary protoliths are found from one layer to the next; OR
ranges mostly between >20 and >50 [OR = molecular 2Fe2O3
× 100/(2Fe2O3 + FeO)]. These OR variations are uncorrelated
with fO2 and do not indicate that large, order-of-magnitude gradients
in fO2 and redox state existed or were preserved between layers during
metamorphism. The other 6% of the samples contain ilmenite, lack
magnetite, and had low fO2 0–1 order of magnitude below QFM
in the stability field of graphite. They are characterized by combinations of the following: large fluid HF/H2O; metasomatic,
tourmaline-bearing veins; absence or rarity of primary organic
matter; and crosscutting late metamorphic shear zones rich in
carbonaceous material. Such observations suggest that locally low
fO2 conditions may have been related to the influx of reducing fluids
from elsewhere in the area.
KEY WORDS: metamorphism;
redox; Barrovian; Scotland; oxygen fugacity
∗Corresponding author. Telephone: +1-203-432-3171. Fax: +1-203432-3134. E-mail: [email protected]
†Present address: Division of Earth and Planetary Sciences,
California Institute of Technology, MC 170-25, 1200 E. California
Blvd., Pasadena, CA 91125, USA.
INTRODUCTION
The chemical and mineralogical evolution of the crust
during orogenesis is critically dependent upon the redox
state of metamorphic fluids ( James, 1955; Chinner, 1960;
Buddington & Lindsley, 1964; Miyashiro, 1964; French,
1966; Eugster & Skippen, 1967; Thompson, 1970; Harte,
1975; Ohmoto & Kerrick, 1977; Rumble, 1978; Lamb
& Valley, 1985; Frost, 1991a, 1991b; Brenan et al., 1995;
Harlov et al., 1997; Ague, 1998). The classical Barrovian
type locality in Glen Clova, Angus, Scotland, is a superb
example of redox controls on metamorphism (Chinner,
1960). Based on field relations, petrography, and the
chemical systematics of minerals and bulk rocks, Chinner
(1960) showed that large differences in bulk-rock oxidation ratio (OR) existed at local and regional scales
between metasediments during amphibolite facies metamorphism in Glen Clova [OR = molecular 2Fe2O3 ×
100/(2Fe2O3 + FeO)]. These differences were attributed
to chemical contrasts inherited, at least in part, from the
sedimentary protoliths, and were taken as evidence that
significant differences in fO2 and fluid composition existed
between layers during metamorphism (Chinner, 1960).
Rock oxidation ratio variations in Glen Clova have
been unequivocally established, but important questions
remain regarding the metamorphic fluids. For example,
what were the quantitative variations in the fugacities of
O2 and sulfur species such as S2 in the region? Did large
variations in fO2 and fluid composition exist between
intercalated metasedimentary layers? Could metamorphic fluid–rock interaction between layers have overcome characteristic buffer capacities inherited from
variable protoliths and act to homogenize redox states
of fluids over the field area? In an effort to address these
 Oxford University Press 2001
JOURNAL OF PETROLOGY
VOLUME 42
questions, we present a petrologic study of the Glen
Clova rocks focused on quantification and interpretation
of regional variations in the values of intensive variables
during amphibolite facies Barrovian metamorphism. Our
results suggest that regional mapping of metamorphic
fO2 and fluid compositions is useful for assessing possible
redox state histories for the Barrovian type area.
Throughout the text, we use the term ‘redox state’ in a
general way to refer to the fO2 and composition of
metamorphic pore fluids.
NUMBER 7
JULY 2001
some pre-sillimanite zone kyanite (Chinner, 1960, 1961;
McLellan, 1985).
The oxide mineral assemblages provide a firm basis
for subdivision of rock types. Chinner (1960) described
assemblages of ilmenite + magnetite, ilmenite + magnetite + hematite, and hematite + magnetite. In addition, we recognize a fourth variety containing ilmenite
alone.
PETROGRAPHIC RELATIONS
GEOLOGIC RELATIONS
The metapelitic and metapsammitic rocks of Glen Clova
form part of the Dalradian Supergroup and were deformed and metamorphosed during the Grampian orogeny (Harte et al., 1984; Fettes et al., 1986; Rogers et al.,
1989; Dempster & Bluck, 1995); peak metamorphism
occurred at >470 Ma (Oliver et al., 2000; Baxter et al.,
2000). The timing of prograde porphyroblast growth
relative to the four main deformational events (D1–D4) in
the area can be summarized as follows (McLellan, 1989):
(1) garnet: mostly syn-D2, some syn- to post-D3; (2) staurolite: syn- to post-D2; (3) kyanite: pre- to syn-D3; (4)
sillimanite: mostly post-D3. Some rocks also contain staurolite and kyanite that formed after sillimanite zone
metamorphism and, presumably, after D3 deformation
(see Chinner, 1961; McLellan, 1985). The inferred bedding and the primary planar deformational fabric are
subhorizontal (dips are <>30°) and, thus, the rocks are
part of the ‘Flat Belt’ of the Scottish Highlands (Harte et
al., 1984).
The Glen Clova area contains much loose float; we
made every effort to sample only from fresh, in situ
outcrops. We investigated rocks in and around Chinner’s
(1960) zone of ‘Hematite-bearing gneisses and associated
semipelitic gneisses’ (Fig. 1), which our mapping indicates
is at least 200–250 m thick. Chinner (1960) reported that
at the margins of this zone the hematite-bearing rocks
are intimately interbedded at the inch scale with hematitefree, graphite-bearing rocks, but we note that most of
the graphitic rocks described in his paper are located
2–5·5 km away from the hematite-bearing zone. We were
unable to confirm or refute the presence of such inchscale interbedding during the course of our field and
laboratory work.
McLellan (1985) inferred the main metamorphic reactions for the area. ‘Peak’ sillimanite zone conditions
were >650–700°C at >6 kbar (McLellan, 1985), were
probably accompanied by some degree of anatexis
(McLellan, 1989), and may have required heat input by
fluid advection (McLellan, 1989). No change in grade is
evident across the field area, and most rocks retain
We present new compositional data for silicate and oxide
minerals in 33 rocks from the Glen Clova area. Because
Chinner (1960) provided data for garnet, muscovite,
biotite, and rhombohedral oxide for his sample 11, we
also consider this sample. Symbols are defined in Table
1, mineral assemblages are listed in Table 2, mineral
compositions are given in Tables 3–6, and analytical and
data reduction procedures are described in the Appendix.
Excellent petrographic descriptions of silicate assemblages have been given by many workers (e.g. Harry,
1958; Chinner, 1960, 1961; Harte & Johnson, 1969;
McLellan, 1985, 1989). Consequently, we focus on the
oxides and sulfides.
Prograde phases
The rhombohedral oxide (RHOMOX) minerals range
primarily between ilmenite (Ilm) and hematite (Hem) in
composition. RHOMOX are found throughout the rock
matrix and as inclusions in garnet, aluminosilicate mineral, staurolite, and tourmaline porphyroblasts. Most
crystals are subhedral to euhedral and 10–900 m long.
Unaltered RHOMOX in magnetite-bearing rocks contain exsolution lamellae; in many samples, these lamellae
underwent additional fine-scale exsolution themselves
(Fig. 2). The multiple sets of exsolution features almost
certainly developed during progressive cooling from
‘peak’ T, as expected from phase relations in the Fe2O3–
FeTiO3 system (e.g. Burton, 1991), and suggest, although
do not prove, that little chemical alteration of the grains
occurred during cooling (Harlov et al., 1997). Rocks
without magnetite contain Ilm (XIlm >0·95) that lacks
exsolution features.
Cubic solid solutions vary primarily between the magnetite (Mag) and ulvöspinel (Usp) endmembers, but are
dominated by Mag (XMag > >0·95). Mag crystals are
found in the rock matrix and, more rarely, as inclusions
within porphyroblasts. Most grains are subhedral to
euhedral and 10 and 1100 m in diameter. Rare martite
alteration of probable supergene origin is present in a
few samples.
Rutile (Rt) is uncommon. Crystals are generally subhedral, >100 m long, and found in rock matrices
1302
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Fig. 1. Location map. Unpatterned area, amphibolite facies metasedimentary rocks; cross-hatch pattern, post-peak metamorphic intrusions
(‘Newer’ igneous rocks); stipple, para-amphibolite (‘Green Beds’). Samples used for northwestern, central, and southeastern profiles enclosed by
dashed lines. Post-metamorphic, shallow-level, porphyritic dikes and sills omitted for clarity.
and as inclusions within porphyroblasts. Rt crystals are
typically rimmed or partially replaced by RHOMOX. It
is unclear if these rim or replacement textures developed
during prograde metamorphism or if they are retrograde.
Additional Rt parageneses that are almost certainly retrograde are described below.
Some samples contain small amounts (<>1 vol. %)
of sulfides. Pyrite (Py), pyrrhotite (Po), and chalcopyrite
(Ccp) crystals are found mainly in the rock matrix and
much less commonly as inclusions within porphyroblasts.
Crystals range from anhedral to euhedral and most are
between 5 and 1000 m in diameter. Ccp, Po, and Nisulfide are also found as tiny rounded ‘droplets’ within
some Py. Evidence for some low-temperature supergene
alteration of sulfides to Hem and limonite is present in
most samples. Because supergene reactions may have
modified the prograde sulfide assemblages (e.g. alteration
of Po to Py), interpretation of metamorphic sulfide assemblages is somewhat hazardous.
Table 1: List of symbols
Symbol
Definition
Symbol
Definition
Alm
Almandine
Pyr
Pyrophanite
Adr
Andradite
Qtz
Quartz
Ann
Annite
RHOMOX
Rhombohedral oxide
Bt
Biotite
Rt
Rutile
Ccp
Chalcopyrite
Sil
Sillimanite
Gk
Geikielite
Sps
Spessartine
Grs
Grossular
St
Staurolite
Grt
Garnet
Usp
Ulvöspinel
Hem
Hematite
Zrn
Zircon
Ilm
Ilmenite
Ky
Kyanite
a
Activity
Mag
Magnetite
f
Fugacity
Ms
Muscovite
P
Pressure
Phl
Phlogopite
T
Temperature
Po
Pyrrhotite
X
Mole fraction
Prp
Pyrope
Py
Pyrite
OR
Rock oxidation ratio
Retrogression of Fe–Ti oxide phases
All mineral symbols except Gk, Pyr, and RHOMOX from Kretz
(1983).
Small, Ilm-rich (XIlm > >0·85) RHOMOX ± Rt grains
are commonly found (1) along cleavages and margins of
biotite and, more rarely, muscovite crystals, and (2) in
contact with and as partial rims around Mag grains. This
Ilm ± Rt probably formed from retrograde Ti release
1303
JOURNAL OF PETROLOGY
VOLUME 42
Grt
Ky
Sil
RHOMOX
Mag
Py
36A
X
X
X
X
37C-2
X
X
X
X
X
X
38A
X
X
X
X
X
X
39B
X
X
X
X
X
40A
X
X
X
X
41C
X
X
X
X
X
42A
X
X
X
X
X
X
45A
X
X
X
X
Po
X
X
X
X
?
X
47A-1
X
X
X
X
48E
X
X
X
X
X
49A
X
X
X
50B
X
X
X
51B
X
X
X
X
X
52C
X
X
54A
55A
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
57B
X
X
X
X
X
58A
X
X
X
X
X
X
59A
X
60A
X
61B
X
62A
X
X
X
X
X
X
X
X
X
X
X
X
X
X
63A-2
X
X
X
X
64B-2
X
X
X
65A
X
X
66A
X
X
67A
X
X
68A
X
X
69B
X
X
70A
X
X
71A
X
X
X
Chin11∗ X
X
X
PRESSURE AND TEMPERATURE
X
X
X
X
X
X
?
X
X
X
X
X
X
46B
56A
Ccp
X
X
JULY 2001
of Rt, Ti-rich Ilm, Ti-poor Hem, and, rarely, titanite.
These pseudomorphs are inferred to have originally been
prograde crystals because the pseudomorphic replacement has generally preserved relict exsolution textures characteristic of cooling from high T. Fe2O3–FeTiO3
phase relations predict that at T below >500°C, Tipoor Hem and Ti-rich Ilm should coexist (e.g. Burton,
1991). Thus, the pseudomorphed RHOMOX probably
comprise retrograde minerals that may have equilibrated
to some degree with nearby retrograde Ilm ± Rt associated with micas and Mag. RHOMOX grains that
are far from micas and Mag or that are preserved as
inclusions within garnet, aluminosilicate, or tourmaline
porphyroblasts are typically not pseudomorphed.
Table 2: Prograde mineral assemblages
Sample
NUMBER 7
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
n.r.
X
X
X
X
n.r.
n.r.
All samples contain quartz, plagioclase, biotite, and muscovite. Sulfide assemblages not listed if strong limonite alteration present. n.r., not reported; ?, possible trace.
∗Sample 11 of Chinner (1960).
from micas and Mag (e.g. Ague & Brimhall, 1988). The
small Usp contents of the Mag grains probably reflect,
at least in part, such retrograde Ti loss. It is unclear if
retrograde Ilm nucleated and grew during cooling, or if
pre-existing prograde metamorphic grains were compositionally modified during cooling.
Nearby matrix RHOMOX are largely pseudomorphed
by fine-grained assemblages containing variable amounts
P–T estimates for selected samples were computed with
the TWEEQU program of Berman (1991) using the
program’s default thermodynamic data and activity models. Mole fractions were computed using version 0·95 of
the CMP program available at the TWEEQU Web site
(www.gis.nrcan.gc.ca/twq.html). The mole fraction of
hydroxyl in micas was adjusted to include F and Cl
substitution for OH (e.g. Ague & Brimhall, 1988; McMullin et al., 1991). We calculated simultaneous solutions of:
(I) the garnet–biotite Fe–Mg exchange (GARB; Alm +
Phl = Py + Ann) and the SGAM reactions of McMullin
et al. (1991; Alm + Ms = Ann + 2 Sil + Qtz, Prp +
Ms = Phl + 2 Sil + Qtz); or (II) GARB and GRAIL
(Alm + 3 Rt = Sil + 2 Qtz + 3 Ilm). Method (II) was
applied to sample 62A, which contains rutile that appears
not to be retrogressed (see above). Reactions involving
grossular in garnet were not used because grossular
contents are generally small and subject to significant
uncertainties stemming from estimated andradite contents.
Results cluster in the sillimanite field (Fig. 3), averaging
657°C and 5·7 kbar; we used 660°C and 5·7 kbar to
calculate fluid composition herein. T was also estimated
using GARB at 5·7 kbar for five samples that lack aluminosilicate minerals (36A, 45A, 49A, 50B, 52C). The
average T of 668°C agrees with the simultaneous P–T
estimates of Fig. 3. Because larger garnet grains preserve
bell-shaped Mn growth zoning profiles, the temperature–
time evolution of these rocks was insufficient to cause
complete diffusional equilibration of cations in garnet.
Our P–T results are fully consistent with those of McLellan (1985).
REACTIONS
A fundamental conclusion of Chinner (1960) was that,
for any given sample, silicates and oxides reacted together
1304
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Table 3: Rhombohedral oxide compositions
Sample
Type∗
Ti
Al
Cr
Fe3+
V
Fe2+
Mn
Mg
Zn
Total
XIlm
XHem
XGk
XPyr
36A
37C-2
38A
39B
40A
40A
41C
42A
45A
46B
47A-1
47A-1
48E
49A
50B
51B
52C
54A
55A
56A
57B
57B
58A
58A
59A
60A
61B
61B
61B
62A
63A-2
64B-2
65A
66A
66A
66A
66A
66A
67A
68A
68A
68A
69B
70A
71A
71A
Chin11†
M
GS
M
M
IG
RM
IG
IG
M
IG
IG
RM
IG
M
IG
IG
M
IGT
IGK
IK
M
M
M
RM
IG
IG
IGi
IG
ISM
IG
IG
IG
IG
IGi
IG
ISM
IK
IK
M
M
IG
PKSt
IK
IK
IG
ISM
—
0·735
1·730
1·292
1·791
1·298
1·753
1·016
0·596
0·883
1·940
0·915
1·941
0·643
0·800
0·628
1·015
1·592
0·802
0·803
0·866
0·835
1·378
0·970
1·734
0·869
0·738
0·822
1·682
1·782
1·957
1·007
0·911
0·932
1·395
1·711
1·701
1·738
1·661
0·761
1·097
1·325
1·829
0·622
0·922
1·726
1·820
0·516
0·005
0·001
0·002
0·002
0·006
0·002
0·002
0·010
0·010
0·007
0·006
0·002
0·009
0·005
0·003
0·002
0·001
0·005
0·007
0·009
0·004
0·001
0·004
tr.
0·002
0·002
0·002
0·001
0·001
tr.
0·000
0·004
0·006
0·001
0·001
0·001
0·002
0·003
0·004
0·005
0·001
tr.
tr.
0·006
0·001
tr.
n.d.
0·007
0·004
0·006
0·005
n.d.
n.d.
0·002
0·004
0·003
0·001
n.d.
n.d.
0·008
tr.
0·002
0·002
0·004
0·004
0·003
0·004
0·002
0·002
0·010
0·006
0·002
0·005
0·019
0·002
0·002
0·001
0·001
0·001
0·006
0·002
0·002
0·004
0·002
0·002
0·002
0·004
0·006
0·005
0·004
0·004
0·002
0·002
n.d.
2·505
0·522
1·395
0·405
1·397
0·493
1·958
2·786
2·211
0·112
2·164
0·115
2·686
2·385
2·734
1·956
0·805
2·377
2·375
2·247
2·312
1·235
2·034
0·521
2·252
2·507
2·321
0·633
0·431
0·083
1·977
2·167
2·115
1·201
0·572
0·589
0·518
0·669
2·465
1·787
1·334
0·331
2·740
2·136
0·544
0·358
2·968
0·012
0·009
0·014
0·005
n.d.
n.d.
0·006
0·007
0·010
tr.
n.d.
n.d.
0·012
0·008
0·005
0·009
0·008
0·010
0·009
0·008
0·012
0·007
0·012
0·006
0·005
0·010
0·013
0·001
0·004
0·002
0·004
0·005
0·008
0·006
0·002
0·005
0·003
0·004
0·006
0·010
0·010
0·006
0·012
0·010
0·002
0·001
n.d.
0·715
1·692
1·278
1·737
1·257
1·736
0·996
0·573
0·864
1·855
0·878
1·867
0·625
0·686
0·585
0·997
1·555
0·791
0·794
0·852
0·826
1·359
0·960
1·706
0·846
0·720
0·795
1·629
1·767
1·904
0·956
0·886
0·922
1·354
1·644
1·670
1·703
1·614
0·716
1·089
1·295
1·796
0·602
0·906
1·702
1·787
0·516
0·018
0·030
0·007
0·043
0·020
0·010
0·017
0·019
0·012
0·080
0·032
0·073
0·009
0·112
0·041
0·017
0·035
0·009
0·008
0·013
0·008
0·016
0·008
0·023
0·019
0·017
0·026
0·018
0·008
0·050
0·048
0·023
0·009
0·039
0·033
0·025
0·017
0·016
0·042
0·006
0·026
0·025
0·012
0·015
0·011
0·018
n.d.
0·001
0·009
0·005
0·010
0·022
0·006
0·001
0·002
0·004
0·004
0·005
0·002
0·007
0·002
tr.
tr.
0·001
0·001
tr.
0·001
0·001
0·002
0·001
0·002
0·003
0·001
tr.
0·032
0·006
0·002
0·003
0·002
0·001
0·002
0·032
0·005
0·016
0·030
0·002
tr.
0·002
0·005
tr.
tr.
0·012
0·014
n.d.
0·002
0·001
0·001
0·002
n.d.
n.d.
0·002
0·002
0·003
0·001
n.d.
n.d.
0·002
0·001
0·002
0·002
0·001
0·002
0·001
0·001
0·001
0·001
0·001
0·003
0·002
tr.
0·001
0·003
0·001
tr.
tr.
tr.
tr.
0·001
0·002
0·001
0·002
0·001
0·001
0·002
0·002
0·003
0·008
0·001
0·001
tr.
n.d.
100·00
99·26
99·97
99·11
99·30
98·99
99·43
101·48
98·84
99·53
99·03
97·38
99·30
100·58
100·33
99·99
100·00
100·97
101·18
100·84
101·12
100·81
100·00
99·92
100·28
99·73
100·29
100·68
100·10
100·35
100·05
99·75
99·85
100·40
99·87
100·12
99·69
98·80
100·65
100·12
99·73
100·07
101·09
100·12
100·24
100·34
98·59
0·360
0·849
0·643
0·872
0·630
0·868
0·500
0·289
0·436
0·930
0·439
0·934
0·316
0·345
0·294
0·501
0·780
0·398
0·399
0·428
0·415
0·682
0·484
0·857
0·425
0·362
0·401
0·817
0·885
0·953
0·479
0·445
0·464
0·679
0·824
0·837
0·854
0·809
0·360
0·548
0·651
0·902
0·306
0·456
0·852
0·894
0·258
0·630
0·131
0·351
0·102
0·350
0·123
0·491
0·700
0·556
0·028
0·542
0·029
0·676
0·598
0·685
0·491
0·202
0·597
0·597
0·565
0·581
0·309
0·512
0·131
0·564
0·629
0·585
0·158
0·108
0·021
0·495
0·543
0·531
0·301
0·143
0·148
0·130
0·168
0·618
0·449
0·335
0·083
0·688
0·537
0·136
0·090
0·742
tr.
0·005
0·003
0·005
0·011
0·003
tr.
0·001
0·002
0·002
0·003
0·001
0·004
0·001
tr.
tr.
tr.
tr.
tr.
tr.
tr.
0·001
tr.
0·001
0·002
tr.
tr.
0·016
0·003
0·001
0·002
0·001
tr.
0·001
0·016
0·002
0·008
0·015
0·001
tr.
0·001
0·003
tr.
tr.
0·006
0·007
n.d.
0·009
0·015
0·004
0·022
0·010
0·005
0·009
0·010
0·006
0·040
0·016
0·036
0·004
0·056
0·020
0·008
0·018
0·005
0·004
0·006
0·004
0·008
0·004
0·011
0·009
0·008
0·013
0·009
0·004
0·025
0·024
0·011
0·005
0·019
0·017
0·013
0·009
0·008
0·021
0·003
0·013
0·013
0·006
0·008
0·006
0·009
n.d.
Six oxygen formula basis. n.d., not determined; tr., trace (<0·001). Fe2+, Fe3+ calculated based on stoichiometry. Total is total
weight percent incorporating calculated FeO and Fe2O3.
∗GS, at garnet–sillimanite contact; IG, inclusions in garnet rim; IGi; inclusions in garnet interior; IGK, inclusions in garnet
and kyanite; IGT, inclusions in garnet and tourmaline; IK, inclusions in kyanite; ISM, inclusions in sillimanite–muscovite
aggregates; M, matrix grain; PKSt; at contact of plagioclase and post-sillimanite kyanite and staurolite; RM, retrograde grain
in matrix.
†From Chinner (1960).
1305
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 7
JULY 2001
Table 4: Magnetite (Mag) and rutile (Rt) compositions
Sample
Phase
Type∗
Ti
Al
Cr
Fe3+
V
Fe2+
Mn
Mg
Zn
62A
Rt
IG
1·954
tr.
0·002
0·053
0·002
—
0·002
tr.
tr.
99·69
37C-2
Mag
M
0·030
0·017
0·002
1·919
0·003
1·026
tr.
0·003
0·001
99·71
0·959
39B
Mag
M
0·003
0·023
0·004
1·963
0·005
1·002
0·001
0·001
tr.
100·47
0·981
40A
Mag
IG
0·045
0·012
n.d.
1·898
n.d.
1·042
0·002
0·001
n.d.
100·29
0·949
40A
Mag
M
0·041
0·005
n.d.
1·914
n.d.
1·040
tr.
0·001
n.d.
99·66
0·957
41C
Mag
M
0·004
0·008
0·002
1·974
0·007
0·994
0·010
tr.
0·001
99·60
0·987
52C
Mag
M
0·005
0·011
0·001
1·971
0·006
1·004
tr.
tr.
0·001
101·20
0·986
54A
Mag
M
0·002
0·008
0·001
1·983
0·004
1·002
tr.
tr.
tr.
100·39
0·992
56A
Mag
M
0·002
0·009
0·001
1·981
0·004
1·002
tr.
tr.
tr.
100·09
0·991
61B
Mag
M
tr.
0·006
0·001
1·987
0·005
1·000
tr.
tr.
0·001
100·36
0·994
66A
Mag
M
0·001
0·005
0·001
1·989
0·004
0·999
tr.
tr.
0·001
100·30
0·994
71A
Mag
IG
0·001
0·018
0·003
1·970
0·006
1·001
tr.
tr.
tr.
100·19
0·985
71A
Mag
M
0·003
0·017
0·001
1·974
0·003
1·003
tr.
tr.
tr.
100·91
0·987
Total
XMag
—
Four oxygen formula basis. n.d., not determined; tr., trace (<0·001). Magnetite Fe2+, Fe3+ calculated based on stoichiometry.
All Fe as Fe3+ for rutile. Total is total weight percent. Total incorporates calculated FeO and Fe2O3 in magnetite.
∗IG, inclusions in garnet; M, matrix grains.
as part of a thermodynamic system. The primary reaction
that we used to estimate fO2 involves both silicates and
magnetite
4Qtz + 2Mag + 2Sil = O2 + 2Alm.
(1)
The main advantages of this reaction are that (1) nonideal
activity corrections for Alm and Mag are well constrained
for the compositional ranges of interest, and (2) Qtz and
aluminosilicates are essentially pure phases. For rocks
lacking aluminosilicate minerals or garnet we used
3Qtz + 2Mag + Ms = O2 + Ann + Alm
(2)
2O2 + 3Py = Mag + 3S2
for Py-bearing assemblages, and
2O2 + 3Po = Mag + 1·5S2
8Qtz + 6Hem + 4Sil = 3O2 + 4Alm
(4)
6Qtz + 6Hem + 2Ms = 3O2 + 2Ann + 2Alm (5)
4Qtz + 6Hem + 4Ms = 3O2 + 4Ann + 4Sil. (6)
However, these reactions are complicated by uncertain
activity–composition relations for RHOMOX (see
below). This uncertainty is particularly problematic for
the Mag–Hem reaction (6Hem = 4Mag + O2) because
fO2 depends on the sixth power of the Hem activity. Thus,
we do not consider this reaction further.
Given an estimate of the fO2, fS2 was estimated using
(8)
for Po-bearing ones. For samples that lack Mag and Py,
we used the reaction 3O2 + 4Po = 2Hem + 2S2.
The fHF /fH2O estimates used measured biotite compositions (Table 5) following Zhu & Sverjensky (1992).
The log10(a Mg2+/a Fe2+) and log10(a Mn2+/a Fe2+) for the fluid
were estimated with
Fe2+ + 1/3Prp = Mg2+ + 1/3Alm
(9)
Fe2+ + 1/3Sps = Mn2+ + 1/3Alm.
(10)
and
2Qtz + 2Mag + 2Ms = O2 + 2Ann + 2Sil. (3)
We also considered the hematite analogs of reactions
(1)–(3):
(7)
Calculation of log10K for reaction (9) involves a slight
extrapolation of fluid species thermodynamic data beyond
the recommended maximum P of 5 kbar (Shock et al.,
1997), but we emphasize that we are concerned with
regional gradients in fluid composition that are independent of log K, not absolute values. Because thermodynamic data for Sps are lacking, we plot log10(a Mn2+/
a Fe2+) −log10K for reaction (10).
ACTIVITY–COMPOSITION
RELATIONS AND THERMODYNAMIC
DATA
We used the activity model of Ghiorso (1990) for RHOMOX. However, Evans & Scaillet (1997) showed that
1306
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Table 5: Biotite (Bt) and muscovite (Ms) compositions
Sample
Phase
Si
Ti
Aliv
Alvi
FeT
Mn
Mg
Ba
Na
K
F
Cl
Total
36A
Bt
2·768
0·121
1·232
0·357
1·218
0·010
1·163
0·008
0·028
0·850
0·060
0·003
96·12
45A
Bt
2·767
0·133
1·233
0·321
1·288
0·018
1·105
0·007
0·027
0·878
0·082
0·001
95·76
46B
Bt
2·751
0·118
1·249
0·381
1·313
0·008
1·049
0·009
0·049
0·829
0·108
0·003
95·94
49A
Bt
2·771
0·128
1·230
0·340
1·289
0·035
1·064
0·008
0·016
0·892
0·068
0·001
95·87
50B
Bt
2·759
0·135
1·242
0·330
1·228
0·008
1·155
0·008
0·031
0·883
0·092
tr.
95·48
52C
Bt
2·756
0·131
1·244
0·337
1·392
0·009
0·993
0·006
0·027
0·883
0·068
0·001
95·61
54A
Bt
2·733
0·107
1·267
0·370
1·247
0·004
1·166
0·004
0·054
0·832
0·064
tr.
96·11
55A
Bt
2·741
0·121
1·259
0·378
1·182
0·004
1·185
0·006
0·040
0·850
0·081
tr.
95·68
56A
Bt
2·735
0·120
1·265
0·381
1·212
0·006
1·150
0·006
0·038
0·855
0·052
tr.
96·11
57B
Bt
2·736
0·151
1·264
0·339
1·292
0·004
1·062
0·006
0·031
0·883
0·096
tr.
95·96
61B
Bt
2·742
0·097
1·258
0·388
1·265
0·002
1·133
0·006
0·042
0·853
0·075
0·005
95·48
62A
Bt
2·756
0·098
1·244
0·392
1·232
0·003
1·148
0·005
0·031
0·868
0·173
0·003
95·88
66A
Bt
2·732
0·135
1·268
0·367
1·246
0·003
1·100
0·009
0·047
0·864
0·094
tr.
95·80
70A
Bt
2·737
0·126
1·263
0·369
1·185
0·004
1·186
0·006
0·043
0·848
0·074
0·001
95·27
71A
Bt
2·730
0·118
1·270
0·416
1·337
0·001
1·008
0·005
0·046
0·804
0·077
0·001
95·98
Chin11∗
Bt
2·696
0·060
1·304
0·558
0·879
0·015
1·378
n.d.
0·076
0·769
tr.
n.d.
95·13
36A
Ms
3·100
0·055
0·900
1·727
0·192
0·001
0·074
0·009
0·125
0·824
0·022
tr.
94·65
45A
Ms
3·108
0·048
0·892
1·709
0·216
0·001
0·085
0·006
0·094
0·864
0·011
tr.
94·19
46B
Ms
3·102
0·060
0·898
1·769
0·120
tr.
0·082
0·005
0·110
0·828
0·024
tr.
94·35
49A
Ms
3·155
0·042
0·845
1·666
0·234
0·001
0·120
0·014
0·067
0·876
0·019
0·001
94·89
50B
Ms
3·122
0·039
0·878
1·730
0·188
0·001
0·090
0·008
0·116
0·839
0·024
tr.
94·52
52C
Ms
3·127
0·040
0·873
1·729
0·192
tr.
0·085
0·009
0·124
0·830
0·017
tr.
94·00
54A
Ms
3·107
0·041
0·893
1·758
0·163
0·001
0·076
0·009
0·183
0·776
0·011
tr.
95·23
55A
Ms
3·097
0·043
0·903
1·743
0·183
0·001
0·081
0·009
0·146
0·809
0·026
tr.
95·10
56A
Ms
3·093
0·045
0·907
1·730
0·199
0·001
0·085
0·010
0·143
0·804
0·011
tr.
94·58
57B
Ms
3·098
0·052
0·902
1·724
0·190
0·001
0·082
0·010
0·127
0·833
0·011
tr.
94·84
61B
Ms
3·109
0·039
0·891
1·745
0·189
tr.
0·080
0·006
0·155
0·795
0·009
tr.
94·73
62A
Ms
3·096
0·055
0·904
1·790
0·101
0·001
0·076
0·008
0·136
0·808
0·009
tr.
94·85
66A
Ms
3·095
0·043
0·905
1·747
0·183
tr.
0·079
0·010
0·133
0·817
0·013
tr.
94·30
70A
Ms
3·108
0·045
0·892
1·719
0·208
0·001
0·088
0·007
0·139
0·808
0·021
tr.
94·95
71A
Ms
3·104
0·051
0·896
1·750
0·167
tr.
0·077
0·008
0·129
0·810
0·011
0·001
94·39
Chin11∗
Ms
3·077
0·028
0·923
1·738
0·205
0·014
0·117
n.d.
0·160
0·765
n.d.
n.d.
94·34
Eleven oxygen formula basis. n.d., not determined; tr., trace (<0·001). All Fe as Fe2+. Total is total weight percent. Oxygen
equivalents of the fluorine and chlorine atoms have not been subtracted from the weight percent totals.
∗Sample 11 of Chinner (1960).
current RHOMOX models overestimate fO2 for highly
oxidized Mt Pinatubo dacites (see below).
Retrograde Ti loss commonly alters the bulk composition of Mag. However, because ‘peak’ metamorphic
T was < >700°C, the expected XMag for spinel coexisting
with the measured RHOMOX compositions would have
been large and between about 0·80 and 1·0 (Andersen
& Lindsley, 1988; Ghiorso & Sack, 1991). The Ti-poor
Mag inclusions in garnet (Table 4) support this idea
because they should have been shielded from retrograde
equilibration with matrix phases and thus preserve prograde compositions. For T of 650–700°C in the Fe3O4–
Fe2TiO4 system, the appropriate Mag activities for our
samples are between about 0·8 and 1·0 (Andersen &
Lindsley, 1988; Ghiorso & Sack, 1991); we used a very
conservative range between 0·75 and 1·0 herein. Other
components such as MgAl2O4 could also have affected
Mag activity, but they are present in only trace amounts
(Table 4); their concentrations are generally significant
only at much higher T (Ghiorso & Sack, 1991).
Po that coexists with Py in the Fe–S system deviates
from ideal FeS by as much as Fe0·875S at high T. However,
natural high-T Po compositions commonly equilibrate
down to very low T, at which many complex phase
1307
0·010
0·000
0·000
0·004
0·007
0·000
3·005
2·988
3·000
2·994
2·961
3·001
2·992
3·022
2·996
2·982
3·000
2·989
2·986
3·005
2·998
2·995
2·990
3·004
2·993
2·990
2·990
3·005
3·002
2·996
2·993
36A
37C-2
38A
39B
40A
41C
42A
45A
46B
47A1
48E
49A
50B
51B
52C
54A
55A
57B
58A
59A
60A
61B
62A
63A-2
64B-2
1308
0·006
0·007
0·000
2·991
2·994
2·993
3·001
2·991
3·041
67A
68A
69B
70A
71A
Chin11†
3·000
3·000
3·041
3·001
3·000
3·000
1·959
1·976
1·902
1·975
1·963
1·965
1·945
1·910
1·931
0·001
0·007
tr.
0·001
tr.
tr.
0·001
tr.
0·002
0·001
0·001
0·001
tr.
tr.
tr.
tr.
tr.
tr.
tr.
0·001
0·001
tr.
tr.
0·001
0·001
tr.
0·001
tr.
tr.
tr.
tr.
tr.
0·001
tr.
Ti
0·023
0·055
0·025
0·036
0·035
0·055
0·090
0·069
0·038
0·020
0·031
0·036
0·039
0·029
0·031
0·042
0·038
0·067
0·036
0·046
0·033
0·027
0·045
0·034
0·076
0·065
0·033
0·049
0·017
0·053
0·022
0·041
0·063
0·018
Fe3+
2·324
1·615
2·117
2·222
2·165
1·953
1·936
2·139
2·189
2·140
2·237
2·249
2·172
2·323
2·162
2·142
2·045
2·068
2·156
2·067
1·888
2·180
2·236
2·199
2·218
2·134
2·098
1·955
2·314
2·108
2·192
1·841
2·039
2·241
Fe2+
0·129
0·427
0·345
0·219
0·314
0·312
0·421
0·195
0·163
0·274
0·176
0·075
0·273
0·106
0·268
0·298
0·441
0·347
0·289
0·326
0·584
0·288
0·206
0·201
0·052
0·311
0·353
0·462
0·096
0·312
0·263
0·561
0·306
0·199
Mn
0·394
0·786
0·423
0·390
0·412
0·472
0·228
0·412
0·412
0·406
0·425
0·396
0·401
0·391
0·425
0·415
0·403
0·384
0·415
0·423
0·297
0·410
0·388
0·416
0·479
0·410
0·398
0·334
0·406
0·438
0·408
0·353
0·422
0·413
Mg
0·158
0·141
0·114
0·170
0·112
0·267
0·394
0·267
0·222
0·182
0·166
0·275
0·159
0·169
0·138
0·148
0·116
0·201
0·142
0·188
0·220
0·128
0·169
0·186
0·270
0·143
0·155
0·204
0·187
0·151
0·135
0·250
0·240
0·138
Ca
3·004
2·969
2·998
3·002
3·003
3·004
2·979
3·013
2·985
3·001
3·003
2·995
3·005
2·990
2·993
3·003
3·005
3·001
3·003
3·005
2·989
3·006
2·999
3·002
3·019
2·997
3·004
2·955
3·002
3·009
2·999
3·005
3·006
2·990
viii
99·99
99·52
99·79
100·18
100·03
100·09
99·95
100·29
99·82
99·99
99·97
100·16
99·87
99·83
100·07
99·91
99·89
100·31
100·39
99·83
99·65
100·33
100·12
99·83
99·75
99·69
100·10
99·13
99·61
100·72
99·82
99·97
99·90
100·00
Total
0·774
0·562
0·706
0·740
0·721
0·650
0·650
0·710
0·733
0·713
0·745
0·751
0·723
0·777
0·722
0·713
0·681
0·689
0·718
0·688
0·631
0·725
0·745
0·732
0·735
0·712
0·698
0·662
0·771
0·701
0·731
0·613
0·678
0·749
XAlm
0·131
0·120
0·141
0·130
0·137
0·157
0·076
0·137
0·138
0·135
0·142
0·132
0·134
0·131
0·142
0·138
0·134
0·128
0·138
0·141
0·099
0·136
0·129
0·139
0·159
0·137
0·132
0·113
0·135
0·146
0·136
0·117
0·140
0·138
XPrp
0·043
0·272
0·115
0·073
0·105
0·104
0·141
0·065
0·054
0·091
0·058
0·025
0·091
0·036
0·089
0·099
0·147
0·116
0·096
0·109
0·195
0·096
0·069
0·067
0·017
0·104
0·117
0·156
0·032
0·104
0·088
0·187
0·102
0·066
XSps
0·040
0·019
0·026
0·038
0·020
0·061
0·087
0·054
0·054
0·050
0·039
0·074
0·033
0·042
0·031
0·028
0·019
0·034
0·029
0·039
0·057
0·029
0·034
0·044
0·051
0·015
0·035
0·044
0·054
0·024
0·034
0·063
0·048
0·037
XGrs
0·012
0·027
0·012
0·019
0·018
0·028
0·045
0·034
0·020
0·011
0·016
0·018
0·019
0·015
0·016
0·021
0·019
0·034
0·018
0·024
0·017
0·013
0·022
0·018
0·038
0·033
0·017
0·025
0·009
0·026
0·011
0·020
0·032
0·009
XAdr
Twelve oxygen formula basis. tr., trace (<0·001). Fe2+, Fe3+ calculated based on stoichiometry. Total is total weight percent incorporating calculated FeO and Fe2O3.
∗Garnet interior.
†Sample 11 of Chinner (1960). Mole fractions from Chinner (1960).
0·009
0·000
0·009
3·007
3·010
3·000
1·979
1·968
1·964
1·961
1·971
1·969
1·958
1·962
1·933
1·964
1·953
1·967
1·973
1·955
1·965
1·923
1·935
1·966
1·951
1·983
1·948
1·978
1·960
1·936
1·982
Alvi
NUMBER 7
0·000
0·025
3·007
3·010
2·975
65A
3·000
3·000
3·002
3·000
3·005
3·004
3·000
3·000
3·000
3·000
3·000
3·005
3·000
3·000
3·000
3·000
3·001
3·000
3·022
3·000
3·000
3·000
3·000
3·000
3·005
iv
VOLUME 42
66A Int.∗
66A
0·000
0·007
0·010
0·002
0·005
0·010
0·000
0·012
0·000
0·006
0·039
0·000
0·008
0·000
0·004
0·018
0·000
0·011
0·014
0·000
Si
Sample
Aliv
Table 6: Garnet compositions
JOURNAL OF PETROLOGY
JULY 2001
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Fig. 3. Pressure–temperature estimates. Ε, method (I); Χ, method
(II) (see text). Samples 46B, 54A, 55A, 57B, 61B, 62A, 66A, 70A, and
71A. Al2SiO5 phase relations from Berman (1988).
Fig. 2. Backscattered electron images of rhombohedral oxides (RHOMOX) included within interior (a) and rim (b) of garnet in sample
66A. RHOMOX inclusions are more ilmenite rich towards rim of
garnet. Ilmenite-rich lamellae, dark gray; hematite-rich lamellae, light
gray. Multiple sets of exsolution lamellae in (a) should be noted.
Composition of hematite-rich domains in (a) (including fine ilmeniterich lamellae): XIlm = 0·32; XHem = 0·67; XGk = 3·9 × 10−4; XPyr =
7·8 × 10−3. Ilmenite-rich domains in (a) (including fine hematite-rich
lamellae): XIlm = 0·84; XHem = 0·13; XGk = 1·0 × 10−3; XPyr = 2·3
× 10−2. White, equant grain at left in (b) is zircon (Zrn).
transitions occur. In view of these retrograde complications, we make the simplification of setting the
activity of Po to 1·0. This simplification introduces some
error, but the activity corrections for even highly nonstoichiometric Po have minor impact on the resolution
of the large variations in fugacities of interest here. Py is
relatively pure and its activity was set to 1·0.
Activity coefficients for garnet components, Ann, and
Ms were computed following Berman (1990), McMullin
et al. (1991) and Chatterjee & Froese (1975), respectively.
Because the rocks are in the sillimanite zone and
nearly all samples contain Sil, thermodynamic data for
sillimanite were used in the calculations; use of Ky has
little impact on the results. Reconnaissance analyses
indicate that the aluminosilicate minerals are relatively
pure Al2SiO5 (XAl2SiO5 >0·99) and thus can be modeled
using unit activity. Qtz was assumed to be pure SiO2.
Thermodynamic data for Mag and most silicate components were taken from Berman (1988). Data for Hem,
Ann, Alm, and F-bearing micas were taken from Ghiorso
(1990), McMullin et al. (1991), Berman (1990), and Zhu
& Sverjensky (1992), respectively, and are internally
consistent with Berman (1988). Results computed with
these standard state properties are nearly identical to
those computed using recently revised properties for Alm
and Mag (Berman & Aranovich, 1996). For example,
the two datasets yield fO2 values for reaction (1), the main
reaction we used to estimate fO2, that differ by only 0·04
log units. We used data from Johnson et al. (1992), as
updated by Shock et al. (1997) and Sverjensky et al. (1997),
for sulfide minerals, graphite, and most fluid species.
Data for the gases CO, HF, and COS are not included
in the Johnson et al. (1992) database and were thus taken
from Robie et al. (1979), Stull & Prophet (1971), and
Wagman et al. (1982), respectively.
Speciation calculations were carried out for C–O–H–S
fluids composed of H2O, H2, CO2, CO, CH4, COS, S2,
SO2, and H2S. Fugacity and activity coefficients for H2O
and CO2 were from Kerrick & Jacobs (1981), whereas
the other fugacities were computed using the equations
of Shi & Saxena (1992) assuming ideal mixing of nonideal gas species. Solution of the non-linear system of
speciation equations was carried out using standard procedures (e.g. Ferry & Baumgartner, 1987; Holloway,
1987).
FUGACITIES AND MINERAL
CHEMISTRY
Oxygen fugacity
Results are given in terms of QFM defined here as the
difference, at 660°C and 5·7 kbar, between the log10 fO2
1309
JOURNAL OF PETROLOGY
VOLUME 42
Fig. 4. QFM computed using hematite-bearing reactions (QFM
Hem) vs QFM computed using magnetite-bearing ones (QFM Mag).
Results obtained using reactions (1) and (4) for aluminosilicate +
garnet-bearing assemblages, reactions (2) and (5) for aluminosilicatefree assemblages, and reactions (3) and (6) for garnet-free assemblages.
When possible, QFM Hem was estimated using compositions of
RHOMOX grains included in garnet rims or in aluminosilicate minerals. For samples lacking RHOMOX inclusions in porphyroblasts,
compositions of RHOMOX in the matrix as far removed as possible
from micas were used (Table 3). Length of black bars corresponds to
QFM Mag calculated for the range in magnetite activities from 0·75
to 1·0 (see text). In this and other plots, analytical variability in Grt
almandine content yields uncertainties in QFM Mag of >±0·03
log10 units (not shown). QFM Hem and QFM Mag agree reasonably
well for samples with RHOMOX compositions on the ilmenite-rich
side of the ilmenite–hematite solvus (37C-2, 39B, 52C, 61B, 66A, 71A).
High fO2 sample 11 of Chinner (1960) denoted by open rectangle.
Samples 46B and 62A lack magnetite, but are plotted on the 1:1
correlation line for comparison purposes.
for a sample and that for the quartz–fayalite–magnetite
(QFM) buffer (QFM log10 fO2= −17·46). Because T–
fO2 trajectories for rocks are generally subparallel to that
for QFM, use of QFM notation largely removes uncertainties caused by errors in T estimates and, therefore,
highlights real fO2 variations between samples. Furthermore, the reactions used to estimate fO2 are fairly
insensitive to P variations, thus minimizing errors caused
by errors in the P estimate.
QFM values computed using Hem-bearing reactions
(QFM Hem) are plotted against the corresponding
Mag-bearing reactions (QFM Mag) in Fig. 4. Most
rocks have QFM Hem between +1·75 and +3·0 and
QFM Mag of +1·8 to +2·3. However, QFM values
for the Mag-free samples 46B and 62A are 2–3 orders
of magnitude lower than those for the other rocks (QFM
−0·25 to −0·75; Fig. 4).
Closer inspection reveals that the QFM Hem estimates are as much as >10 times greater than the
NUMBER 7
JULY 2001
QFM Mag estimates. The six samples that do cluster
near the 1:1 correlation line (37C-2, 39B, 52C, 61B,
66A, 71A) all have RHOMOX compositions that lie on
the Ilm-rich side of the Hem–Ilm solvus. Agreement
between QFM Hem and QFM Mag for these samples
may be due to the fact that most experiments in the
Fe2O3–FeTiO3 system have used Ilm-rich compositions
(see Ghiorso, 1990). Consequently, the activity–
composition relations for Ilm-rich RHOMOX are probably reasonably well understood. We note that QFM
Hem for sample 39B, which plots somewhat farther from
the 1:1 line than the other five samples, is based on the
composition of matrix RHOMOX. The RHOMOX in
this sample may have undergone some retrograde Ti
gain by exchange with matrix phases during cooling
(see above), and its QFM Hem may be somewhat
underestimated.
The samples that plot well above the 1:1 line have
compositions that are either on the solvus or to the Hemrich side (Fig. 4). Problems with current RHOMOX
activity models near the solvus have been documented
(Evans & Scaillet, 1997), and the Hem-rich side of
the solvus remains to be fully explored experimentally.
Consequently, the QFM Hem values for these samples
are inferred to be less reliable than those for Ilm-rich
compositions. Furthermore, the RHOMOX in samples
that plot well above the 1:1 line contain complex exsolution textures. It is possible that these grains originally
straddled the solvus and were thus composed of two
coexisting RHOMOX compositions at peak T. In this
case, our ‘integrated’ compositions, which are averaged
over all exsolution domains in RHOMOX grains, would
be in error and could not be used in thermodynamic
calculations. Resolution of the above issues requires a
better understanding of phase relations among rhombohedral oxides.
To avoid problems with the thermodynamic treatment
of RHOMOX, we focus on QFM values computed
using Mag-bearing reactions (Table 7). The two exceptions are samples 46B and 62A, which lack Mag
and require QFM estimates based on Hem-bearing
reactions. Because these two samples contain Ilm-rich
RHOMOX that lacks exsolution features, they should
be treated reasonably well with the activity model and
standard state thermodynamic data used herein.
The agreement between QFM estimates made using
reactions (1), (2), and (3) for Mag-bearing samples, or
reactions (4), (5), and (6) for Mag-free samples (Table 7),
indicates that the thermodynamic data and activity models employed are internally consistent.
Mineral chemistry
Prograde RHOMOX vary from XHem >0·02 to >0·75
(Table 3). XHem values for inferred retrograde RHOMOX
1310
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Table 7: Fugacity and rock compositions
Sample
QFM(1)
36A
QFM(2)
QFM(3)
FeO
Fe2O3
OR
7·2
3·00
31·6
2·05–2·30
HF/H2O
Elevation (m)
0·100
325
37C-2
1·86–2·11
38A
1·80–2·05
505
39B
1·85–2·10
405
40A
1·86–2·11
365
41C
1·91–2·16
445
42A
1·96–2·21
45A
46B
460
1·93–2·18
−0·36∗
47A1
1·92–2·17
48E
1·84–2·09
440
−0·34†
−0·33‡
7·3
3·49
0·266
263
0·417
315
35·4
370
450
49A
1·96–2·21
0·194
390
50B
1·94–2·19
0·294
310
1·70–1·95
0·227
430
0·136
480
51B
1·78–2·03
52C
340
54A
1·89–2·14
1·87–2·12
1·84–2·09
55A
2·01–2·26
1·97–2·22
1·93–2·18
4·6
2·84
43·5
0·230
510
1·82–2·07
8·0
4·41
39·7
0·047
385
0·355
420
56A
57B
1·87–2·12
58A
1·91–2·16
59A
2·02–2·27
60A
1·87–2·12
61B
1·69–1·94
62A
−0·75∗
1·77–2·02
1·67–1·92
460
490
1·73–1·98
−0·69†
1·77–2·02
−0·63‡
6·6
3·77
40·9
10·2
2·76
21·7
0·216
490
525
8·0
1·81
18·5
0·595
450
5·6
4·28
51·1
7·6
2·85
28·9
63A-2
1·91–2·16
64B-2
1·80–2·05
410
65A
1·85–2·10
66A
1·94–2·19
67A
2·17–2·42
450
68A
1·81–2·06
485
69B
1·87–2·12
70A
1·92–2·17
1·91–2·16
1·90–2·15
5·3
3·52
46·0
0·190
445
71A
1·70–1·95
1·68–1·93
1·66–1·91
5·2
2·23
32·4
0·279
385
Chin11§
2·51–2·76
2·56–2·81
2·60–2·85
3·7
12·01
74·8
350
370
1·87–2·12
1·79–2·04
0·326
400
495
QFM (1), QFM (2), QFM (3) computed using reactions (1), (2), and (3), respectively, except where noted otherwise. Range
in values corresponds to the magnetite activity range 0·75–1·0. FeO and Fe2O3 are bulk-rock weight percents. OR, rock
oxidation ratio. HF/H2O is the difference in log10( fHF /fH2O) between the sample and log10( fHF /fH2O) for a standard biotite having
XMg = 0·5 and log10(XF/XOH) = −1·5 at 660°C and 5·7 kbar.
∗Computed using reaction (4).
†Computed using reaction (5).
‡Computed using reaction (6).
§Sample 11 of Chinner (1960).
intergrown with micas are smaller than >0·15 and are
significantly smaller than the XHem of crystals included in
nearby porphyroblasts (Table 3).
‘Felted mats’ of Sil intergrown with micas and, in some
cases, retrograde St or Ky commonly contain significant
quantities of RHOMOX. These yield XHem values that
are equal to or smaller than those for inclusions in
porphyroblasts (Table 3). It is unclear if the smaller
1311
JOURNAL OF PETROLOGY
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NUMBER 7
JULY 2001
Fig. 6. Spessartine/almandine mole fraction ratio (XSps/XAlm) vs
QFM. High fO2 sample 11 of Chinner (1960) denoted by open
rectangle.
Fig. 5. Chemical traverses from core to rim across typical garnet
in sample 66A. VI and VIII denote cation sums for six-fold and
eight-fold coordinated sites, respectively.
values resulted from input of reducing fluid, retrogression,
or both. Retrograde effects are difficult to rule out because
the RHOMOX are closely intergrown with micas.
Samples 57B and 68A appear to have two coexisting
prograde RHOMOX compositions that occur as discrete grains. Using the solvus diagram of Ghiorso
(1990), compositions in 68A suggest a T near 680°C,
in good agreement with the T estimates based on
thermobarometry, whereas compositions in 57B are
inconsistent with equilibration at a single T. The
significance of these results is unclear given continuing
uncertainties regarding T–activity–composition relations
for RHOMOX.
Garnet Fe3+ is generally largest in cores and smallest
in rims (Fig. 5). Fe3+ and Ti4+ decrease with Ca (Fig.
5), such that the ratio of grossular to andradite
component remains fairly constant. Garnet Fe3+ contents are small and no correlation between QFM
and garnet Fe3+ is evident, consistent with Chinner
(1960). The Fe3+ systematics are apparently sensitive
mostly to crystal chemical factors and the availability
of Ca, rather than fO2.
Garnet Mn content is in general positively correlated
with rock Fe3+/Fe2+ (Chinner, 1960), although not
necessarily with fO2. The XSps/XAlm ratio is small for
QFM <0, whereas a wide range of XSps/XAlm values
seem compatible with QFM of >+2 (Fig. 6). Sample
11 of Chinner (1960) is somewhat anomalous relative
to our samples and has the largest XSps/XAlm and fO2
(Fig. 6). We note in this regard the possibility that
the bulk garnet composition given by Chinner (1960)
may be more Mn rich than the peak-T composition
at garnet rims.
The XMg/XFeT ratio for biotite changes little over the
QFM range between −1·0 and +2·3 (Fig. 7; FeT
is all Fe as FeO). However, XMg/XFeT increases at the
largest QFM, consistent with the expectation that
higher fO2 should favor more Mg-rich compositions
(e.g. Wones & Eugster, 1965). The XMg/XFeT of muscovite
(Ms) decreases between QFM −1·0 and +2·3, and
then increases somewhat for higher QFM (Fig. 7).
Fe3+ is the dominant form of Fe in Glen Clova Ms
(Chinner, 1960). Under reducing conditions, the total
Fe content of Ms is relatively small because the amount
of Fe3+ available is limited relative to more oxidized
conditions. Consequently, XMg/XFeT ratios for Ms in
the low-fO2 rocks are large. The slight increase in XMg/
XFeT above QFM +2 reflects the stability of magnesian
compositions at higher fO2.
Changes in RHOMOX inclusion
compositions from core-to-rim in garnet
Some samples were examined in detail to determine
if RHOMOX inclusion compositions changed from
core to rim in garnets. This effort was hampered by
the relatively small size and paucity of inclusions in
many of the garnet crystals. None the less, RHOMOX
included in the interiors of garnet porphyroblasts in
samples 61B and 66A were measured to have significantly smaller XIlm than RHOMOX inclusions near
1312
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Fig. 8. Rhombohedral oxide (RHOMOX) ilmenite mole fractions
for samples 61B and 66A. Χ, RHOMOX inclusions in garnet; Α,
inclusions in kyanite rims; Φ, inclusions in sillimanite + muscovite
aggregates. Generalized changes in RHOMOX compositions during
garnet growth denoted by dashed arrows.
orders of magnitude less than the other rocks (Fig. 9;
see below).
Fig. 7. Mg/FeT (all Fe as Fe2+) mole fraction ratio vs QFM for
biotite (Bt; upper panel) and muscovite (Ms; lower panel). High
fO2 sample 11 of Chinner (1960) denoted by open rectangles.
garnet rims, or those included within Sil + Ms
aggregates and Ky rims (Figs 2 and 8; Table 3). One
interpretation of this shift towards larger XIlm is that
reducing fluids, either locally or externally derived,
caused reduction of the mineral assemblage during
garnet growth. Unfortunately, uncertainties in activity–
composition relations for RHOMOX preclude full
quantification of possible fO2 changes recorded by the
increases in XIlm, although calculations using the Ghiorso
(1990) model suggest that drops in fO2 as large as an
order of magnitude occurred during garnet growth.
Because RHOMOX compositions between the small
XIlm values in Grt interiors and the large XIlm values
in rims were not observed, any such reduction was
probably rapid compared with the rate of Grt growth.
Relationships between fO2 and fS2
Most sulfide-bearing samples record fS2 at or slightly
below the Py–Po buffer (Fig. 9). The average fS2 of
Py-bearing rocks is slightly higher than that for Pobearing ones, consistent with thermodynamic predictions. The two samples with the smallest QFM
(46B, 62A) are the only ones that could have been in
equilibrium with graphite. These record fS2 values 2–3
REGIONAL VARIATIONS
Regional variations in fO2
Regional fO2 profiles were constructed for the northwest,
central, and southeast parts of the field area (Fig. 1).
Because the fabric of the rocks is nearly horizontal,
sample data were projected onto vertical sections such
that the profiles are essentially perpendicular to the fabric.
A small number of samples (46–49) are located between
the northwest and central profile areas and are thus not
shown in Figs 10 and 11.
QFM for the northwestern and southeastern profiles
shows no large variations as a function of elevation (Fig.
10a and c). The main feature of the central profile is the
strong localized minimum in QFM associated with
sample 62A (Fig. 10b).
Regional variations in activities of Mg, Fe,
and Mn species in solution
Surprisingly, the calculated log10(a Mg2+/a Fe2+) of the fluid
varies little and shows no significant correlations with
elevation or QFM (Fig. 11). Nearly all values are within
±10% of the mean. Other related ratios, such as log10
(a MgCl+/a FeCl+), would show exactly the same patterns of
regional variation [of course, the absolute values would
be different from log10(a Mg2+/a Fe2+)]. These results strongly
suggest that no large gradients in fluid Mg/Fe existed
between layers. QFM for all samples must have been
low enough so that Fe2+ species were dominant, otherwise
1313
JOURNAL OF PETROLOGY
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NUMBER 7
JULY 2001
Fig. 9. PyPo vs QFM. PyPo is the difference between log10
fS2 for the sample and log10 fS2 for the pyrite–pyrrhotite buffer at
660°C and 5·7 kbar. For rocks that contain magnetite (all those at
QFM above graphite stability), gray and black bars denote pyritebearing and pyrrhotite-bearing samples, respectively. Length of bars
corresponds to QFM calculated for the range in magnetite activities
from 0·75 to 1·0 (see text). It should be noted that only the two
magnetite-free, pyrrhotite-bearing samples (46B, 62A) fall within the
stability field of graphite.
the log10(a Mg2+/a Fe2+) would vary more strongly as a
function of fluid Fe3+/Fe2+.
In contrast, log10(a Mn2+/a Fe2+) varies significantly and
shows a weak positive correlation with fO2 (Fig. 11). In
most cases the log10(a Mn2+/a Fe2+) variations are not abrupt
and restricted to individual beds, but rather are gradational at the 10–100 m scale. Examples include the
gradual decrease in log10(a Mn2+/a Fe2+) from 460 to 360 m
in the northwest profile (Fig. 11a), and the ‘C’-shaped
variation between 300 and 425 m in the southeast profile
(Fig. 11c). The sharper discontinuity in the central profile
corresponds to the extremely low fO2 sample 62A, and
may be of metasomatic origin (see below). Metamorphic
fluid–rock interaction was insufficient to homogenize
fluid Mn2+/Fe2+ in the sequence, possibly owing to low
Mn concentrations in fluids (i.e. large solid–fluid partition
coefficient for Mn).
DISCUSSION OF f O 2 SYSTEMATICS
Chinner (1960) used the rock oxidation ratio (OR) as an
index of rock oxidation state, where OR = 2Fe2O3 ×
100/(2Fe2O3 + FeO) (molecular basis). Although clearly
powerful, OR is not a direct indicator of fO2 because
rocks of differing OR can equilibrate at the same fO2.
For example, for equilibrium at a given P and T, pore
fluid in a rock that contains 99% Hem and 1% Mag
Fig. 10. Regional profiles of QFM across layering for the northwestern, central, and southeastern parts of the field area (Fig. 1). Z is
the elevation of the sample in meters above sea level. Length of black
bars corresponds to QFM calculated for the range in magnetite
activities from 0·75 to 1·0 (see text, Table 7). Sample 62A plots offscale at QFM −0·75. Samples 61B and 66A discussed in text.
would be at the same fO2 as fluid in a rock containing
99% Mag and 1% Hem, even though the OR values for
the two rocks differ considerably.
In fact, a wide range of OR values, between about 20
and 50, are compatible with essentially the same QFM
of about +2 (Fig. 12). Consequently, the observation of
intimately interbedded rocks of differing OR (Chinner,
1960) does not necessarily indicate that the rocks had
very different fO2 and fluid compositions. Extremes of the
OR range are, however, associated with extremes of
fO2. Sample 62A, equilibrated at QFM <0, has OR
<20, whereas sample 11 of Chinner (1960) has an OR
of >75 and QFM > +2·5. These extremes of the
fO2 range are discussed in a later section.
The homogeneity of the bulk of the metamorphic fO2
values can be interpreted in two ways. One possibility is
that the lack of large variations reflects a protolith that
was fairly homogeneous with respect to mineral and fluid
compositions to start with (although OR values must
1314
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
The second possibility is that the sequence was originally more heterogeneous, such that sharp, layer-tolayer variations in redox state controlled by protolith
composition existed. Metamorphic fluid–rock interaction
between layers must then have acted to smooth these
short wavelength variations so that the entire sequence
attained a relatively uniform redox state. The observed
increases in the XIlm and decreases in the XHem of RHOMOX inclusions from core to rim in garnet (Figs 2 and
8) suggest that some syn-metamorphic changes in redox
state did occur.
If redox states were homogenized during metamorphism, the reaction histories would have been critically dependent on fluid compositions and the amount
of fluid transport. The highly oxidized, QFM +2 fluids
of Mag-bearing rocks would have been fairly inefficient
oxidizers or reducers because such fluids consist mostly
of H2O and CO2, and are poor in many species needed
for redox reactions including H2, CH4, and CO (Wood
& Walther, 1986; Frost, 1991a; Ague, 1998). For example,
the reactions
3H2 + 3Hem + Ms + 3Qtz = Ann +
Alm + 3H2O
(11)
0·75CH4 + 3Hem + Ms + 3Qtz = Ann + Alm +
0·75CO2 + 1·5H2O
(12)
Fig. 11. Elevation (Z) in meters above sea level vs fluid log10(a Mn2+/
a Fe2+) − log10 K [reaction (10)] and log10(a Mg2+/a Fe2+) [reaction (9)] for
northwestern, central, and southeastern profiles.
require either 1 mole of H2 or 0·25 moles of CH4 to
reduce 1 mole of Hem and produce garnet and biotite.
Typical QFM +2 Glen Clova fluids, however, probably
had XH2 of only >10−4 and vanishingly small XCH4 <
>10−7 (Table 8). Consequently, large fluid–rock ratios
would have been needed to homogenize the rock package.
Quantification of these ratios is difficult owing to the
uncertain activity–composition relations for RHOMOX,
but calculations using simple model reactions such as (11)
and (12), and representative fluid compositions (Table 8)
and modes (Chinner, 1960) imply that if homogenization
occurred, minimum fluid–rock ratios were >102 (Wood
& Walther, 1986; Ague, 1998). Fluid exchange between
beds would probably have been dominated by diffusion
and mechanical dispersion across layers (e.g. Ague &
Rye, 1999; Baxter & DePaolo, 2000), although advection
may also have played a role. Modeling results suggest that
disruption of protolith bed-scale fugacity heterogeneities
may occur in as little as 102−103 years (Ague, 1998).
Fig. 12. Rock oxidation ratio (OR) vs QFM. High fO2 sample 11 of
Chinner (1960) denoted by open rectangle.
have differed from bed to bed). Metamorphism of such
a sequence would not be expected to produce large
gradients in fO2 and fluid composition between layers,
unless ‘exotic’, nonequilibrium fluids from outside the
sequence were introduced.
PETROGENESIS OF SAMPLES AT
THE EXTREMES OF THE f O 2 RANGE
Rocks equilibrated at extremely low fO2
Two samples (46B, 62A), making up >6% of the sample
set, are from layers that equilibrated at fO2 2–3 orders
of magnitude lower than the rest of the rocks. The layers
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JOURNAL OF PETROLOGY
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NUMBER 7
JULY 2001
Table 8: Fluid mole fractions
Species
Representative
Representative
Representative
Representative
Sample 62A
QFM +1·8
QFM +1·8
QFM +2·4
QFM +2·4
QFM −0·75
XH O = 0·50∗
XH O = 0·95∗
XH O = 0·50∗
XH O = 0·95∗
1·18×10−4
1·78×10−4
5·93×10−5
8·91×10−5
3·01×10−2
−1
−1
−1
−1
8·04×10−1
−2
1·71×10−1
−5
2·10×10−3
−9
1·53×10−2
−3
2·68×10−3
−4
7·55×10−9
−5
1·33×10−4
−6
2
H2
H2O
CO2
CO
CH4
H2S
SO2
COS
5·00×10
−1
4·96×10
−4
2·13×10
−7
1·28×10
−3
3·39×10
−5
6·00×10
−4
3·02×10
2
9·50×10
5·00×10
−2
−1
4·44×10
4·98×10
−5
−4
4·35×10
1·07×10
−8
−9
5·90×10
8·09×10
−3
−3
5·10×10
1·70×10
−5
−4
6·00×10
2·39×10
−5
−4
6·17×10
1·52×10
−6
9·50×10
4·71×10
2·29×10
3·91×10
2·55×10
2·39×10
3·25×10
1·56×10
1·56×10
1·56×10
1·56×10
3·11×10−9
aGraphite
5·39×10−3
1·10×10−3
1·36×10−3
2·90×10−4
1·00×100
−0·25
−6
2
S2
PyPo†
−6
2
−0·25
−0·25
−0·25
−2·95
∗Calculated for representative range in XH2O based on McLellan (1985).
†PyPo is the difference between log10fS2 for the fluid and log10fS2 for the pyrite–pyrrhotite buffer at 660°C and 5·7 kbar.
are at least meter scale in thickness, but whether contacts
with surrounding rocks are sharp, gradational, tectonic,
or otherwise could not be determined in the field. Either
the sedimentary or diagenetic precursors to these two
samples were highly reduced, or reduction resulted from
metamorphic fluid–rock interaction. Some evidence from
graphite and fluorine systematics suggests the latter possibility.
Graphite, inferred to have been derived from sedimentary organic matter, is common in low-OR rocks far
removed (kilometer scale) from the study area (Chinner,
1960). In contrast, graphite is absent from the lowest
fO2 sample (62A). This sample probably originally lacked
or contained only traces of organic matter and, thus, its
protolith may not have been highly ‘reduced’. Consequently, the low fO2 could have been produced during
metamorphism.
Traces of graphite may be included within a few mica
grains in 46B, but the identification is uncertain. On the
other hand, the sample is cut by millimeter- to centimeterwide, post-peak metamorphic shear zones containing
abundant carbonaceous material, plagioclase, muscovite,
and chlorite, and minor calcite (Fig. 13). Reducing fluids
were therefore able to infiltrate the rock, albeit not
necessarily on the prograde path.
The low fO2 samples record larger HF/H2O than the
others (Fig. 14). In addition, sample 46B is cut by veins
that contain abundant F-rich tourmaline crystals as much
as 1 cm long of probable metasomatic origin (see Dutrow
et al., 1999). We suggest that the HF/H2O systematics
reflect the local infiltration of chemically active fluids
into the low fO2 rocks, and draw the connection that these
fluids also drove syn-metamorphic reduction. The HF/
H2O relations might also have been inherited from sedimentary protoliths, but it is difficult to envision what
sedimentary process could produce a negative correlation
between HF/H2O and QFM (Fig. 14).
The above evidence suggests that infiltration of reduced
metamorphic fluids drove reaction and lowered fO2 in
samples 46B and 62A. If these fluids were at or near
graphite saturation, they would have contained abundant
H2 and CH4 (sample 62A, Table 8). Because these species
are powerful reducing agents, alteration would have
occurred even if fluid–rock ratios were rather small—of
the order of one or less (Wood & Walther, 1986; Ague,
1998). The fluid could have been derived from graphitic
rocks, such as those that surround the sequence studied
here, or from reduced magmas that intruded during peak
metamorphism. If the low fO2 was produced during metamorphism, the localized reduction suggests channelized
fluid infiltration.
High fO2 sample
Sample 11 of Chinner (1960) has both large OR and
QFM relative to our samples (Fig. 12). These characteristics were probably inherited from the protolith
because (1) the large Mn and Fe contents of the rock
suggest a protolith with high bulk-rock Fe2O3/FeO [see
discussion by Chinner (1960)], and (2) no local or regional
sources for highly oxidized metamorphic fluids needed
1316
AGUE et al.
HIGH fO2 DURING SILLIMANITE ZONE METAMORPHISM
Fig. 13. Photomicrograph of post-peak metamorphic shear zone in sample 46B. Deformed muscovite grains (Ms) lie within chlorite-rich matrix.
Note abundant dust-like carbonaceous material. Field of view 1·35 mm wide; plane-polarized light.
Fig. 14. HF/H2O vs QFM. HF/H2O is the difference in
log10( f HF /f H2O) between the sample and log10( f HF /f H2O) for a standard
biotite having XMg = 0·5 and log10(XF/XOH)= −1·5 at 660°C and
5·7 kbar. High fO2 sample 11 of Chinner (1960) denoted by open
rectangle. F in this sample was below detection (Chinner, 1960). Its F
content was set to the smallest value measured from our sample suite
and thus probably represents a maximum.
to oxidize the rock are evident. One possibility is that
sample 11 is an unusual relic that escaped to some degree
redox equilibration with the rest of the sequence, perhaps
because of extremely low porosity and/or permeability.
Another possibility is that the bulk garnet composition
of Chinner (1960) used in our calculations has lower
Fe/Mn and almandine activity than the peak-T rim
composition. If so, then the fO2 estimate made using the
bulk garnet composition would be somewhat too high.
CONCLUSION
Estimated fO2 for the bulk of the study area was about
+2 log10 units above the quartz–fayalite–magnetite buffer
(QFM >+2) at sillimanite zone conditions, although
rare low fO2 rocks (QFM <0) were also present. The
high fO2 of QFM >+2 is comparable with the
highest known from arc settings (e.g. dacite from Mt
Pinatubo; Evans & Scaillet, 1997). With the exception
of the rare low fO2 rocks, variations in metamorphic
fO2 across the study area were generally small, about
±0·3 log10 units. This minimal variability may reflect
either (1) a protolith that was initially homogeneous with
respect to redox state or (2) a protolith characterized by
significant differences in redox state from bed to bed
that was homogenized by metamorphic fluid–rock
interaction. Rhombohedral oxide (RHOMOX) inclusions in garnet grains that become richer in ilmenite
and poorer in hematite towards garnet rims suggest
that some syn-metamorphic reduction of rock occurred.
In detail, the QFM >+2 rocks comprise intimately
intercalated, graphite-free metasedimentary layers with
whole-rock oxidation ratios (OR) that vary mostly
between >20 and >50 from one layer to the next
[OR = molecular 2Fe2O3 × 100/(2Fe2O3 + FeO)].
These OR variations are probably inherited from
sedimentary protoliths, are uncorrelated with fO2, and
do not indicate that large, order-of-magnitude gradients
in fO2 and redox state existed between these layers
during metamorphism. As pointed out by Chinner
(1960), metamorphic fluid–rock interaction did not
homogenize fluid compositions at scales of several
kilometers, because the high fO2, QFM >+2 rocks
are in regional contact with Fe2O3-poor, graphitebearing lithologies. The rare low fO2 rocks (QFM <0)
in our field area that occur among the dominant high
fO2 (QFM >+2) rocks may have been reduced by
regionally migrating fluids derived from these graphitebearing surroundings.
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VOLUME 42
ACKNOWLEDGEMENTS
We thank D. M. Rye, E. L. McLellan, R. L. Masters,
and C. J. Carson for thoughtful discussions, M. S. Ghiorso
for sharing both software and insightful advice for calculating activity–composition relations among Fe–Ti oxides, B. W. Evans, B. R. Frost, and an anonymous referee
for critical reviews, and S. S. Sorensen for editorial
handling. Financial support from the National Science
Foundation, Division of Earth Sciences, grant EAR9405889, is gratefully acknowledged.
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APPENDIX: ANALYTICAL AND DATA
REDUCTION PROCEDURES
Mineral composition determinations and electron microscopy used the JEOL JXA-8600 electron microprobe at
Yale University. Quantitative analyses employed wavelength-dispersive spectrometers, 20 nA Faraday cup current, 15 kV accelerating voltage, natural and synthetic
standards, off-peak and fluorescence-corrected, non-linear mean atomic number background corrections, and
( z) matrix corrections. Electron microscope image
processing and qualitative energy-dispersive chemical
analyses were carried out with systems mounted on the
JXA-8600; the early stages of our study used a Kevex
Delta system and the later stages used an EDAX Phoenix
system.
All Fe was assigned to FeO for the matrix corrections.
This procedure can lead to inaccurate results for Fe3+rich phases if the ‘excess’ oxygen associated with Fe2O3
is not accounted for. Consequently, oxygen was incorporated by difference into the matrix correction iterations for RHOMOX and Mag such that the sum of
all oxides and the ‘excess’ oxygen was 100 wt %. Similarly,
H2O was specified by difference in the matrix corrections
for micas such that the sum of the oxides and the H2O
was 100 wt %.
Bulk compositions for the exsolved RHOMOX were
estimated by making as many as 40 spot determinations
on several grains per thin section. A 5 or 10 m diameter
beam was used in nearly all cases, but a focused beam
NUMBER 7
JULY 2001
was required for several fine-grained samples. Data were
typically collected along linear traverses perpendicular to
the long axes of the exsolution lamellae. At least five spot
analyses on 1–2 grains per thin section were carried out
for unexsolved RHOMOX (10 m beam), Rt (5 m
beam), Mag (focused beam), garnet (focused beam), and
micas (15 or 20 m beam). Linear traverses across garnet
involved as many as 50 determinations. Unless noted
otherwise, reported garnet analyses are for inferred ‘peak’
compositions having the highest Mg and lowest Mn
contents (usually several tens of m from the rim).
Structural formulae for RHOMOX and magnetite
were computed assuming 2 moles of cations per three
oxygens for RHOMOX, and 3 moles of cations per four
oxygens for magnetite. These procedures were judged
successful because total oxide sums incorporating estimated FeO and Fe2O3 are always near 100 wt %
(Tables 3 and 4). FeO and Fe2O3 in garnet can be
estimated in several ways; we assumed 2 moles of octahedrally coordinated cations per 12 oxygens. This
procedure gives the best results when tested against the
wet chemical determinations by Deer et al. (1992), and
gives total cation site occupancies very close to the
ideal number of three for the four-fold and eight-fold
coordinated sites (Table 6).
Bulk-rock X-ray fluorescence analyses were performed
by X-ray Assay Laboratories, Don Mills, Ontario,
Canada. FeO was determined using standard wet chemical titrations.
1320